Click
Here
GEOPHYSICAL RESEARCH LETTERS, VOL. 33, L20502, doi:10.1029/2006GL027484, 2006
for
Full
Article
Permafrost carbon: Stock and decomposability of a globally significant
carbon pool
S. A. Zimov,1 S. P. Davydov,1 G. M. Zimova,1 A. I. Davydova,1 E. A. G. Schuur,2
K. Dutta,2 and F. S. Chapin III3
Received 6 July 2006; revised 11 September 2006; accepted 19 September 2006; published 27 October 2006.
[1] The magnitude of future CO2-induced climate warming
is difficult to predict because of uncertainties in the role of
ecosystems and oceans as CO2 sources and sinks. Siberia has
extensive areas (1 106 km2) of deep (up to 90 m) deposits of
organic-rich frozen loess (wind-blown silt) that accumulated
during the Pleistocene but have not been considered in most
global carbon (C) inventories. Similar deposits occur less
extensively in Alaska. Recent warming at high latitudes
causes this permafrost (permanently frozen ground) to thaw,
raising questions about the fate of C in thawing permafrost.
Here we show that Siberian loess permafrost contains a large
organic C pool (450 GT—more than half the quantity in the
current atmosphere) that decomposes quickly when thawed,
and could act as a positive feedback to climate warming.
Citation: Zimov, S. A., S. P. Davydov, G. M. Zimova, A. I.
Davydova, E. A. G. Schuur, K. Dutta, and F. S. Chapin (2006),
Permafrost carbon: Stock and decomposability of a globally
significant carbon pool, Geophys. Res. Lett., 33, L20502,
doi:10.1029/2006GL027484.
1. Introduction
[2] It is difficult to predict the magnitude of future CO2induced climate warming because of uncertainties in the
role of ecosystems and oceans as CO2 sources and sinks.
Siberia has extensive areas (1 106 km2) of deep (up to
90 m) deposits of organic-rich frozen loess (wind-blown
silt) that accumulated during the Pleistocene (Figure 1)
[Romanovsky, 1993; Schirrmeister et al., 2002; Sher et al.,
2005; Tomirdiaro, 1980; Vasil’chuk and Vasil’chuk, 1998,
1997]. The existence, spatial distribution, and carbon (C)
concentrations of these sediments are well documented in the
Russian literature but there has been only one preliminary
estimate of their C stock [Zimov et al., 1997], so they have
never been considered in global C inventories. Similar
deposits occur less extensively in Alaska [Pewe and
Journaux, 1983]. Estimates of high-latitude soil C stocks
are based largely on studies of the top 1 – 2 m of soils,
predominantly in North America [Batjes, 1996; Gorham,
1991; Melillo et al., 1995; Prentice et al., 2001; Smith et
al., 2004]. Values for these C stocks have then been extrapolated to a pan-arctic basis [Batjes, 1996; Smith et al., 2004].
1
North-East Scientific Station, Pacific Institute for Geography, Far-East
Branch, Russian Academy of Sciences, Cherskii, Russia.
2
Department of Botany, University of Florida, Gainesville, Florida,
USA.
3
Institute of Arctic Biology, University of Alaska, Fairbanks, Alaska,
USA.
Copyright 2006 by the American Geophysical Union.
0094-8276/06/2006GL027484$05.00
Here we provide the first quantitative estimate of C stocks in
Siberian loess permafrost and show that the pool is large
enough (450 GT C—more than half the quantity in the
current atmosphere) to warrant consideration in global C
inventories.
[3] High-latitude soil C stocks are of particular interest
because recent warming thaws permafrost (permanently
frozen ground) [Osterkamp and Romanovsky, 1999;
Romanovsky et al., 2001], raising questions about the fate
of permafrost C [Zimov et al., 2006]. Here we show that the C
contained in Siberian loess permafrost decomposes quickly
when thawed. We describe observations that provide insights
into the processes by which C accumulates in or disappears
from Siberian loess permafrost and discuss the implications
of changes in permafrost C for the global C cycle.
2. Methods
[4] Near Cherskii on the Kolyma River in Northeast
Siberia, we studied the process of C incorporation into
permafrost by comparing six grasslands that spanned a wide
range of depositional environments, soil moisture, and thaw
depths (70 – 150 cm) to understand the past formation of
loess permafrost when grassland ecosystems were widespread. We sampled three soil profiles per community in
September 1998 (the time of maximum thaw) from the soil
surface to about 20 cm beneath the maximum thaw depth.
Communities along a sediment deposition gradient included
an annually flooded Calamagrostis riverbank (meadow 1);
two additional Calamagrostis meadows further from the
river that receive less sediment input (meadows 2 and 3); a
moist herb meadow (meadow 4); a dry herb meadow (dry
meadow); and a steppe at the base of a steep south-facing
slope (dry steppe) [Zimov et al., 1999].
[5] We collected frozen loess permafrost (termed yedoma
in Siberia) that formed during the Pleistocene from deep soil
profiles recently exposed by river or coastal erosion. Sites
were selected along a current climate gradient from a higharctic coastal site to boreal forest. Sites were: Chukochii
Mys (high arctic 1); Zelenyi Mys (forest tundra 1) Duvannyi
Yar (forest tundra 2); and Stanchikovskii Yar (boreal forest).
At the last site, the middle of the profile had a buried
organic horizon from an interstadial. At Chukochii Mys and
Duvannyi Yar, we also sampled sediments beneath former
thermokarst lakes where the Pleistocene ice wedges had
melted during the Holocene (causing the total soil thickness
to decrease due to loss of ice wedge volume from the loess,
forming the lake) (former lakes 1 and 2). The unfrozen loess
sediments beneath the lakes refroze later in the Holocene
after the lakes drained.
L20502
1 of 5
L20502
ZIMOV ET AL.: VULNERABILITY OF PERMAFROST CARBON
Figure 1. The distribution of yedoma loess sediments in
North Siberia (shaded). The study area along the Kolyma
River is highlighted by the black rectangle. Modified from
Romanovsky [1993].
[6] Subsamples from both modern grasslands and yedoma
permafrost were dried and analyzed for C using the WalkleyBlack method [Allison, 1965]. We transported soils to the
U.S. and reanalyzed a set of 153 samples (from this and other
studies) on a Carlo Erba CHN analyzer to develop a regression relationship between the Walkley-Black method, which
measures oxidizable compounds to total organic carbon, as
measured on the CHN analyzer. To be conservative, we used
the Walkley-Black C values, which averaged 13% less than
values from the CHN analyzer, because these values were
available for all our samples and were methodologically
comparable to values obtained from the Russian literature.
Samples included all organic C in the profile, including roots
and buried organic matter; no carbonates were detected by
acidification and subsequent remeasurement with the CHN
analyzer. We present data on a gravimetric dry-mass rather
than volumetric basis to avoid biases associated with regional
or temporal variations in soil ice content.
[7] Immediately after yedoma sample collection in July
1998, we thawed and initiated respiration measurements of
selected subsamples in 1 L (8 L for Duvannyi Yar samples)
cylinders. One set of samples was measured during one
summer at a stable temperature (+4°C) to simulate temperature conditions that thawed soils would experience within
the soil profile; a second set was incubated over 6 seasons at
ambient temperatures (0 – 24°C in summer) to simulate conditions that near-surface soils would experience after permafrost thaw. Incubations naturally warmed and thawed during
the 6 growing seasons (May –Oct, 1998– 2003) and were
frozen during the rest of the year, just as would occur in the
field. Soils were also warmed for 45 days in Nov – Dec 1999,
resulting in 7 warm periods during 6 calendar years. At about
weekly intervals during summer, we placed a lid on the
chamber and measured midday respiration rate with a LICOR 6200 infrared gas analyzer [Zimov et al., 1999]. Water
was added as necessary to maintain soil moisture. We
surface-sterilized another set of frozen yedoma samples
from Zelenyi Mys with 96% alcohol, removed the alcohol
[Gilichinskii et al., 1989], and incubated them at 0°C (n = 3)
and 15°C (n = 3), and measured respiration rate daily to
weekly for one summer.
L20502
[8] We measured in situ soil respiration of recently
thawed unvegetated loess at Duvannyi Yar (two dates)
and Zelenyi Mys (one date) in July, using 0.2 m2 plastic
chambers and a LI-COR 6200 infrared gas analyzer [Zimov
et al., 1999]. We estimated the mass of thawed soil beneath
each chamber (n = 59) from thaw depth (47 ± 6.7 cm) and a
representative bulk density of 1.65 g cm 3 and assumed that
all thawed soil contributed to the observed surface flux.
Bulk density of these mineral soils varied primarily due to
ice crystal content, rather than C content since the latter was
relatively low. Air temperatures ranged from 8– 14°C during these measurements.
[9] A separate incubation was conducted at the University of Florida with yedoma sediments that were collected
and transported frozen. These incubations were conducted
at a constant temperature (15°C) in contrast to the earlier
incubations, and the CO2 respired by microbial activity was
purified and split for 13C analysis using a GasBench attachment for a Finnigan Delta XL Plus mass spectrometer. The
other split was reduced to graphite in a hydrogen atmosphere
with an iron catalyst and sent to the Keck Carbon Cycle
Accelerator Mass Spectrometer facility at the University of
California, Irvine for 14C analysis [Vogel et al., 1987].
Carbon dioxide concentrations were measured with a LICOR 6252, and percent C measured with a Costech elemental
analyzer. Respiration rate is reported as the average rate
observed over the first thirty days, and the isotope values
reflect the integrated CO2 released during that time period.
[10] Our calculations of total organic C in permafrost
include roots and buried organic matter, in contrast to our
earlier calculations [Zimov et al., 1997], which were based on
literature values for soils from which roots had been removed. These sediments currently occupy 1 106 km2
(estimated by digitizing yedoma areas in Figure 1), have an
average thickness of 25 m (10 – 60 m) (estimated from the
literature), an ice-wedge content of about 50% (estimated
from the literature) [Romanovsky, 1993; Schirrmeister et al.,
2002; Sher et al., 2005; Tomirdiaro, 1980; Vasil’chuk and
Vasil’chuk, 1998, 1997], a bulk density of 1.65 g cm 3
(average from our measurements), and a C content of
2.56%, (n = 71 from Siberia and Alaska). At the beginning
of the Holocene, the total C content of these sediments was
about 528 GT. We reduced this estimate to 450 GT for current
yedoma because 50% of the area is occupied by former thaw
lakes with 30% less soil C.
3. Results and Discussion
[11] Grass-dominated tundra-steppe vegetation predominated in Pleistocene landscapes north of 45°N that were free
of glacial ice during cold dry glacial periods [French, 1996;
Velichko, 1973]. A similar climate exists today in northeast
Siberia, where grasslands occur in areas receiving regular
sediment deposition [Zimov et al., 1995]. Yedoma permafrost
forms syngenetically through loess deposition. In other
words, when wind-blown or alluvial materials are deposited
on the soil surface, the bottom of the thawed soil profile,
which includes roots and other organic matter, becomes
incorporated into permafrost. The C concentration in frozen
loess soil beneath the active layer showed relatively little
variation among modern grasslands (1.27– 2.13%) and was
similar between permafrost (1.60 ± 0.07%, n = 30) and the
2 of 5
L20502
ZIMOV ET AL.: VULNERABILITY OF PERMAFROST CARBON
L20502
Figure 2. Vertical profiles of organic C content and respiration rates from Northern Siberian soils.
(a) Modern grassland soils receiving contemporary sediment inputs. Data are means (±SE) of 3 profiles
per site. Arrows denote samples within a profile that are currently in permafrost; all other samples are
within the seasonal active layer. (b) Pleistocene-aged permafrost sediments beneath five profiles that
never thawed since their formation and beneath former thermokarst lakes formed by melting of ice
wedges during the Holocene (causing the soil thickness to decrease). These latter sediments later refroze
after the lakes drained. At Stanchikovskii Yar (boreal forest), the middle of the profile had a high C
content because it was a buried organic soil from an interstadial. There can be many such layers in
yedoma exposures [Schirrmeister et al., 2002; Tomirdiaro, 1980; Vasil’chuk, 2004]. (c) Soil carbon
dioxide flux from laboratory incubations of thawed permafrost sediments and modern grassland soils.
Respiration rate is reported as the average rate observed over the first summer of incubation (average soil
temperature = 4.0°C).
base of the overlying thawed soil (1.66 ± 0.10%, n = 18;
Figure 2a), indicating that soils do not change their C content
as they become incorporated into permafrost in response to
surface mineral deposition.
[12] The average C content of Late Pleistocene permafrost loess was 2.38% ± 0.38 (n = 57) across a climatic
transect from arctic tundra to boreal forest along the Kolyma
River in northeast Siberia (Figure 2b). The average C
content of these sediments was intermediate between measurements from Alaskan yedoma (3.31 ± 0.32%, n = 14) and
previous measurements from northeast Siberia (2.21 ±
0.28%; n = 9) [Zhigotsky, 1982], but lower than in yedoma
Figure 3. Time course of respiration from laboratory incubations of Northern Siberian soils. (a) Late
Pleistocene permafrost sediments from 6 depths at Duvannyi Yar (forest tundra 2; n = 6) incubated at
ambient air temperature (average = 11°C) for six years. Summer temperature is depicted as the blue line at
the top referenced to the right axis. (b) Respiration of Pleistocene loess that had never thawed (Zelenyi Mys,
forest tundra 1, n = 6); sediments beneath former thermokarst lakes (Chukochii Mys, and Duvannyi Yar;
former lake 1 and 2, respectively; n = 3); and soils from modern Calamagrostis meadows (modern
grassland; n = 12).
3 of 5
L20502
ZIMOV ET AL.: VULNERABILITY OF PERMAFROST CARBON
L20502
Figure 4. Laboratory incubation respiration rates and C isotopes of respired CO2 from Northern Siberian soil incubations.
(a) The relationship between respiration rate and bulk soil C content. Data are means (±SE) of 3 samples per site. Sites are a
subset of those denoted in Figure 1, with the addition of a site at Sukharnaya (high arctic 2) and at Medvezhka (high arctic 3).
Three depths from Zelenyi Mys (forest tundra 1) were incubated, including the current surface soil as a reference. (b) The d 13C
and D14C values (left axis) of respired CO2, with corresponding 14C age (right axis). The surface soil had a positive mean
radiocarbon value (modern) and is not shown.
from the Laptev seacoast (mean 5%, minimum 2%, n = 55)
[Schirrmeister et al., 2002]. We also sampled soils beneath
former thermokarst lakes, i.e., lakes that had formed by
thawing ice-rich loess and refroze after the lakes drained.
These soils had profiles that were only half as deep, due to the
loss of ice volume, and had 30% less C than did permanently
frozen loess, due to C loss primarily as Pleistocene-aged
methane and carbon dioxide were emitted from anaerobic
sediments beneath lakes when they thawed in the early
Holocene [Zimov et al., 1997, Walter et al., 2006]. We
calculate the C pool of yedoma permafrost to be 450 GT,
based on our measured C concentration and bulk density and
published reports on the areal extent, thickness, and ice
content of yedoma [Romanovsky, 1993; Tomirdiaro, 1980;
Vasil’chuk, 2004; Vasil’chuk and Vasil’chuk, 1997]. Inclusion of this yedoma C pool stored in deep loess permafrost
approximately doubles previous inventory estimates of soil C
storage in northern latitudes, which typically have
been estimated only for the surface 1 – 2 m [Batjes, 1996;
Gorham, 1991; Melillo et al., 1995; Prentice et al., 2001;
Smith et al., 2004].
[13] Reactivity is just as important as pool size in governing the potential impact of yedoma on the current global C
cycle. We used several independent approaches to document
the respiratory loss of C when frozen yedoma thaws. Each of
these experiments showed that this C was lost quite quickly
when permafrost thawed:
[14] 1. In situ respiration of recently thawed unvegetated
loess averaged 2.3– 5.3 g C m 2 d 1 (3.0– 6.0 mg C kg 1 of
soil d 1) in July, greater than the mid-summer rates currently
observed in soils from productive ecosystems in the region
[Zimov et al., 1999].
[15] 2. In the laboratory, we measured respiration rates at
natural soil temperatures of soils collected from the sites
described above: (a) frozen loess with Pleistocene ice
wedges, (b) refrozen loess from the sediments of a former
thermokarst lake, and (c) modern grassland soils and associated permafrost. All sediments (collected at depths up to
50 m) had high respiration rates (3 – 11 mg C kg 1 of soil d 1;
Figure 2c). The respiration rates of Pleistocene sediments
were generally higher than those of lower horizons of
productive grassland soils. Soil respiration in these incuba-
tions remained high for 6 summer seasons of measurement
(Figures 3a and 3b), indicating that C availability to microbes
remained relatively high throughout the experiment, a
result also observed with modern boreal and tundra soils
of high C lability [Neff and Hooper, 2002; Weintraub and
Schimel, 2003]. The average respiration in the sixth year
was 3.18 mg C kg 1 d 1 (average temperature 14.8°C).
The respiration rates in this experiment include bursts of
respiration associated with initial thaw and annual freezethaw events (as would occur naturally) but are conservative
because we did not measure winter respiration.
[16] 3. Frozen loess that we surface-sterilized to remove
possible contamination by modern microbiota respired at
rates of 2.3 mg C kg 1 d 1 at 0°C and 4.4 mg C kg 1
d 1 at 15°C, indicating rapid initiation of respiratory
activity of live microbes present in Pleistocene-aged permafrost [Gilichinskii et al., 1995].
[17] 4. The unvegetated loess at Zelenyi Mys that had
thawed eight years previously had 39% less soil C (1.10 ±
0.07%, n = 5), averaged across the depth profile, than
adjacent sediments that had remained frozen (1.71 ±
0.20%, n = 3; same bulk density), indicating a rapid and
substantial loss of C after thaw of Pleistocene-aged loess.
The thawed soil had a respiration rate that was 40% less
than that of the permafrost soil (and proportional to remaining soil C content; data not shown), when incubated at the
same temperature.
[18] Another set of incubations, described in more detail
by Dutta et al. [2006], showed that yedoma respiration rate
increased linearly with soil C content, and that the source of
this respiration was ancient C (21000 – 25000 yr BP)
that became available to microbes when the soil thawed
(Figures 4a and 4b).
[19] Each of these experiments showed that Siberian
permafrost C is quite decomposable when thawed, both from
in situ and in laboratory incubations. The large quantities of
readily decomposible C currently present in frozen loess have
important implications for atmospheric CO2 concentrations.
The respiration rates we observed (>3 mg C kg 1 d 1) are
equivalent to 1.1 kg C (T soil) 1 yr 1, assuming the minimum in situ respiration that we observed in recently thawed
yedoma. The average C content that we measured in yedoma
4 of 5
L20502
ZIMOV ET AL.: VULNERABILITY OF PERMAFROST CARBON
soils of Siberia and Alaska was 25.6 kg C T 1 (n = 71). Thus,
the typical release time for C in yedoma could be a few
decades after it thaws if conditions were similar to the
incubations. Substantial permafrost thaw is plausible in
central Siberia, where permafrost temperature (0 to 3°C)
has warmed to near the melting point [Romanovsky et al.,
2001]. Thawing of permafrost can be self-sustaining due to
changes in surface energy balance and the heat released by
soil respiration that add heat and continue the thawing
process [Zimov et al., 1996]. The recent warming [Briffa et
al., 1995] and increased fire frequency [Kasischke et al.,
1999; Vlassova, 2002] that have occurred in Siberia in the
past 30 years could trigger permafrost warming and thaw
similar to recent trends in Alaska [Romanovsky et al., 2002]
and projections for the panarctic [Lawrence and Slater,
2005].
[20] Organic C stored in permafrost has unique properties
that influence its role in the global C budget. The permafrost
C pool accumulates slowly at a rate controlled by sedimentation, can be preserved for hundreds to millions of years
when frozen, and, although deep in the soil profile, is highly
decomposible and can be released quickly when thawed.
These properties suggest that factors inducing high-latitude
climate warming should be mitigated to minimize the risk of a
potentially large CO2 release that would cause a positive
feedback to climate warming.
[21] Acknowledgments. We thank Katey Walter and David McGuire
for critical review of the manuscript and the Russian Fund for Fundamental
Research, the U.S. National Science Foundation, the U.S. Civilian Research
and Development Foundation, the U.S. Department of Energy, and the
National Aeronautics and Space Agency for financial support.
References
Allison, L. E. (1965), Organic carbon, in Methods of Soil Analysis, edited
by C. A. Black, pp. 1367 – 1378, Am. Soc. of Agron., Madison, Wis.
Batjes, N. H. (1996), Total carbon and nitrogen in the soils of the world,
Eur. J. Soil Sci., 47, 151 – 163.
Briffa, K. R., et al. (1995), Unusual twentieth-century summer warmth in a
1,000-year temperature record from Siberia, Nature, 376, 156 – 159.
Dutta, K., et al. (2006), Potential carbon release from permafrost soils of
Northeastern Siberia, Global Change Biology, in press.
French, H. M. (1996), The Periglacial Environment, 2nd ed., Longman,
New York.
Gilichinskii, D. A., et al. (1989), Mikrobiologicheskaya harakteristika pri
izuchenii osadochnyh porod kriolitozony, Izv. Akad. Nauk SSSR Ser.
Geol., 6, 103 – 115.
Gilichinskii, D. A., et al. (1995), Microbial life in permafrost (a historical
review), Permafrost Periglacial Processes, 6, 243 – 250.
Gorham, E. (1991), Northern peatlands: Role in the carbon cycle and
probable responses to climatic warming, Ecol. Appl., 1, 182 – 195.
Kasischke, E. S., et al. (1999), Satellite imagery gives a clear picture of
Russia’s boreal forest fires, Eos Trans. AGU, 80, 141,147.
Lawrence, D. M., and A. G. Slater (2005), A projection of severe nearsurface permafrost degradation during the 21st century, Geophys. Res.
Lett., 32, L24401, doi:10.1029/2005GL025080.
Melillo, J. M., et al. (1995), Global change and its effects on soil organic
carbon stocks, in Role of Nonliving Organic Matter in the Earth’s Carbon
Cycle, edited by R. G. Zepp and C. Sonntag, pp. 175 – 189, John Wiley,
Hoboken, N. J.
Neff, J. C., and D. U. Hooper (2002), Vegetation and climate controls on
potential CO2, DOC and DON production in northern latitude soils,
Global Change Biol., 8, 872 – 884.
L20502
Osterkamp, T. E., and V. E. Romanovsky (1999), Evidence for warming
and thawing of discontinuous permafrost in Alaska, Permafrost Periglacial Processes, 10, 17 – 37.
Pewe, T. L., and A. Journaux (1983), Origin and character of loess-like silt
in unglaciated South-Central Yakutia, Siberia, USSR, U.S. Geol. Surv.
Prof. Pap., N1262, 1 – 46.
Prentice, I. C., et al. (2001), The carbon cycle and atmospheric carbon
dioxide, in Climate Change 2001: The Scientific Basis, edited by J. T.
Houghton et al., pp. 183 – 237, Cambridge Univ. Press, New York.
Romanovsky, N. N. (1993), Osnovy Kriogeneza Litosfery, 336 pp., Izdatelstvo Moscow State Univ., Moscow.
Romanovsky, V. E., et al. (2001), Permafrost temperature dynamics along
the East Siberian transect and an Alaskan transect, Tohoku Geophys. J.,
36, 224 – 229.
Romanovsky, V., et al. (2002), Permafrost temperature records: Indicators
of climate change, Eos Trans. AGU, 83, 589 – 594.
Schirrmeister, L., et al. (2002), Paleoenvironmental and paleoclimatic records from permafrost deposits in the Arctic region of Northern Siberia,
Quat. Int., 89, 97 – 118.
Sher, A. V., S. A. Kuzmina, T. V. Kuznetsova, and L. D. Sulerzhitsky
(2005), New insights into the Weichselian environment and climate of
the East Siberian Arctic, derived from fossil insects, plants, and mammals, Quat. Sci. Rev., 24, 533 – 569.
Smith, L. C., et al. (2004), Siberian peatlands a net carbon sink and global
methane source since the early Holocene, Science, 303, 353 – 356.
Tomirdiaro, S. V. (1980), Lessovo-ledovaia formatsia Vostochnoi Sibiri v
pozdnem pleistotsene i golotsene, 184 pp., Nauka, Moscow.
Vasil’chuk, Y. K. (2004), Yuzhny predel areala povtorno-zhil’nykh l’dov v
Evrazii, Kriosfera Zemli, 3, 34 – 51.
Vasil’chuk, Y. K., and A. C. Vasil’chuk (1997), Radiocarbon dating and
oxygen isotope variations in late Pleistocene syngenetic ice-wedges,
northern Siberia, Permafrost Periglacial Processes, 8, 335 – 345.
Vasil’chuk, Y. K., and A. C. Vasil’chuk (1998), 14C and 18O in Siberian
syngenetic ice-wedge complexes, Radiocarbon, 40, 883 – 893.
Velichko, A. A. (1973), Prirodnyi Process v Pleistocene, 256 pp., Nauka,
Moscow.
Vlassova, T. K. (2002), Human impacts on the tundra-taiga zone dynamics:
The case of the Russian lesotundra, Ambio, 12, 30 – 36.
Vogel, J. S., et al. (1987), 14C background levels in an accelerator mass
spectrometry system, Radiocarbon, 29, 323 – 333.
Walter, K. M., S. A. Zimov, J. P. Chanton, D. Verbyla, and F. S. Chapin III
(2006), Methane bubbling from Siberian thaw lakes as a positive feedback to climate warming, Nature, 443, 71 – 75.
Weintraub, M. N., and J. P. Schimel (2003), Interactions between carbon
and nitrogen mineralization and soil organic matter chemistry in arctic
tundra soils, Ecosystems, 6, 129 – 143.
Zhigotsky, V. Y. (1982), Korennoe izmenenie geokhimii landshaftov na
nizmennostyah Severo-Vostoka SSSR na granice Pleistocene-Holocene,
in Merzlotno-geologicheskie processy i paleogeographiya nizmennostei
Severo-Vostoka Azii, pp. 101 – 110.
Zimov, S. A., et al. (1995), Steppe-tundra transition: An herbivore-driven
biome shift at the end of the Pleistocene, Am. Nat., 146, 765 – 794.
Zimov, S. A., et al. (1996), Siberian CO2 efflux in winter as a CO2 source
and cause of seasonality in atmospheric CO2, Clim. Change, 33, 111 –
120.
Zimov, S. A., et al. (1997), North Siberian lakes: A methane source fueled
by Pleistocene carbon, Science, 277, 800 – 802.
Zimov, S. A., et al. (1999), Contribution of disturbance to high-latitude
amplification of atmospheric CO2, Science, 284, 1973 – 1976.
Zimov, S. A., et al. (2006), Permafrost and the global carbon budget,
Science, 312, 1612 – 1613.
F. S. Chapin, Institute of Arctic Biology, University of Alaska, Fairbanks,
AK 99775-7000, USA.
S. P. Davydov, A. I. Davydova, S. A. Zimov, and G. M. Zimova, NorthEast Scientific Station, Pacific Institute for Geography, Far-East Branch,
Russian Academy of Sciences, Republic Sakha (Yakutia), 678830 Cherskii,
Russia.
K. Dutta and E. A. G. Schuur, Department of Botany, University of
Florida, 220 Bartram Hall, Gainesville, FL 32611, USA. ([email protected])
5 of 5